Transient high strain rate during localised viscous creep in the dry lower
1
continental crust (Lofoten, Norway)
2
Campbell, L.R.*1, Menegon, L1,2. 3
1School of Geography, Earth and Environmental Sciences, Plymouth University, Plymouth PL4 8AA, 4
UK 5
2The Njord Centre, Department of Geosciences, University of Oslo, P.O. Box 1048 Blindern, Norway 6
*Corresponding author: Lucy Campbell ([email protected]) 7
8
Key Points:
9
Viscous creep at high differential stress and strain rate is observed in microstructures within 10
lower crustal mylonitised pseudotachylytes;
11
The high stress and strain rate deformation was transient and localised, as shown by partial 12
development of lower stress microstructures;
13
Pseudotachylytes could support transients in stress and strain rate within strong lower crust, e.g.
14
as observed during postseismic relaxation.
15
Abstract 16
Understanding the ability of the lower crust to support transient changes in stresses and strain rates 17
during the earthquake cycle requires a detailed investigation of the deformation mechanisms and 18
rheology of deep crustal fault rocks. Here we show that lower crustal pseudotachylyte - bearing 19
shear zones are able to accommodate short-term episodes of high strain rate, high stress 20
deformation by accelerated viscous creep, followed by a reduction in stresses to some ambient 21
deformation condition.
22
Quartz microstructure within pseudotachylyte - bearing shear zones in otherwise undeformed 23
granulites from Lofoten, Norway, indicates that dynamic recrystallisation occurred during viscous 24
creep under rapid strain rates and high stresses of ~10-9 s-1 and ~100 MPa, respectively. Lower stress 25
microstructures (i.e. foam-textures) are also recorded in the shear zones, indicating spatial and 26
temporal variations of stress and strain rate during deformation cycles. Both the high and lower 27
stress quartz recrystallization took place under granulite facies conditions of 650-750°C and 0.7-0.8 28
GPa and represent a record of highly localised viscous creep within the lower crust. This implies that 29
lower crustal pseudotachylytes are potentially able to form extremely localised weak zones within 30
strong lower crust, enabling a deep mechanical response to perturbations in stress and strain rate 31
such as those experienced during the seismic cycle, for example seismogenic loading followed by 32
subsequent postseismic relaxation.
33 34
Plain language summary 35
Detailed investigation of the strength and deformation style of fault rocks sourced from the Earth’s 36
lower crust is important to understand how the lower crust reacts to short-term variations in stress 37
and strain rate, which can occur, for example, between earthquakes. Here, we show that solidified 38
pseudotachylytes (initially melts produced due to frictional heating along the fault plane during an 39
earthquake) occurring at depths of 25-30 km in the lower crust can accommodate deformation at 40
particularly high strain rates and high stresses via solid-state creep. We look at pseudotachylytes 41
formed in lower crustal shear zones that are now exhumed in Lofoten, Norway. Deformation 42
microstructures in quartz within these pseudotachylytes have recorded rapid strain rates and high 43
stresses. These microstructures are occasionally transformed into lower stress versions, indicating 44
that during the deformation the stress and strain rate varied through both time and space. Both 45
stages, however, record the same deformation temperatures and pressures, indicating that these 46
are snapshots of ongoing deformation within the lower crust. We conclude that, when the lower 47
crust is strong, pseudotachylytes will form important weak zones that accommodate deformation 48
even during rapid variations in the deformation conditions – for example as occurs during the 49
postseismic period immediately after an earthquake.
50 51
1. Introduction 52
The rheological behaviour of dry lower crustal rocks is commonly characterised by a cyclic interplay 53
between viscous creep (mylonitisation) and brittle, frequently coseismic, fracturing associated with 54
formation of pseudotachylytes (Austrheim, 2013; Hawemann et al., 2018; Jamtveit et al., 2018 55
Menegon et al., 2017; Okudaira et al., 2015; Sibson, 1980; White, 1996; Wex et al., 2019).
56
Earthquakes are effective precursors of ductile shear zones in dry and strong lower crustal regions, 57
as they trigger rheological weakening by grain size reduction via fracturing and comminution 58
alongside potential fluid infiltration (Jamtveit et al., 2019; Petley-Ragan et al., 2019). These 59
processes may facilitate mylonitic creep, dominantly by grain size sensitive deformation localised to 60
the hydrated volume of fractured rocks, commonly consisting of pseudotachylyte veins and their 61
damage zone (Jamtveit et al., 2019; Menegon et al., 2017, Passchier 1982; White 1996). However, 62
mylonitised pseudotachylytes may be overprinted by new generations of pseudotachylytes, thus 63
indicating a cyclical interplay between aseismic creep and coseismic slip along the same structure 64
(Menegon et al., 2017; Wex et al., 2019). This cyclical interplay already implies remarkable 65
oscillations in stress and strain rate during the activity of lower crustal faults and shear zones 66
presumably steered by the earthquake cycle. Whilst this level of cyclical interplay is demonstrated in 67
a number of detailed studies of exhumed lower crustal shear zones (Wex et al., 2019, and references 68
therein), one existing question to address is whether the microstructures of lower crustal fault rocks 69
can preserve a record of the rheological response to transient phenomena away from steady state 70
creep rates. Investigating the microstructure of lower crustal fault rocks resulting from cyclical 71
perturbations in viscous deformation conditions is particularly timely in the light of recent 72
deformation experiments that demonstrate the ability of the recrystallized grain size of quartz to 73
capture transients (Kidder et al., 2016).
74
Transiently elevated strain rates and stresses are observed in creeping faults from both the 75
geological record (e.g. in the form of mutually overprinting pseudotachylytes and mylonites) and, at 76
a more detailed temporal scale, from continuous geodetic observation of crustal deformation. In 77
many cases, geodetic observations of elevated strain rates are ambiguous as to the contributing 78
mechanisms, and could be explained through the inclusion of frictional afterslip in addition to (or 79
instead of) accelerated viscous creep in the mid– to lower-crust (Fagereng & Biggs, 2018; Ingleby &
80
Wright, 2017). It is clearly important, therefore, to understand the controls on where transient 81
changes to crustal deformation rate are feasible, but the deformation mechanisms and materials 82
and that can support rapid strain rates and their subsequent decay are still debated. Important 83
questions are, for example, whether transiently high postseismic strain rates require the presence of 84
a weak material distributed under the seismogenic fault (Ivins, 1996) or if they can alternatively be 85
accommodated within a single lithology with the appropriate rheology (Chopra, 1997), and whether 86
that rheology needs to be linear viscous (Thatcher, 1983), Burgers body (Hearn et al., 2002, Hearn et 87
al., 2009) or power law (Ingleby & Wright, 2017). Therefore, understanding the geological record of 88
transient rheologies in exhumed deep crustal fault rocks promises to be highly useful for 89
constraining fault zone models used to explain surface deformation observations (Bürgmann &
90
Dresen, 2008, Floyd et al. 2016).
91
In this contribution, we utilise an exhumed lower-crustal shear zone network from the eastern 92
Nusfjord region (Lofoten, Norway) to investigate processes that may facilitate rapid strain rates over 93
short time scales within longer-term periods of steady state creep. Additionally we consider the 94
mechanisms of producing and maintaining localisation of viscous deformation within dry, feldspar- 95
rich lower crustal rocks, representative of intracontinental lower crustal shear zones in metastable, 96
impermeable, and mechanically strong granulites (c.f. Austrheim et al., 1996; Jamtveit et al., 2018).
97
We look at the deformation of pseudotachylytes, which are initially produced by coseismic frictional 98
melting but, once present, provide grain size and lithological contrasts that are susceptible to 99
mylonitic overprinting. The Nusfjord pseudotachylytes have already been well-constrained in terms 100
of their mineralogy, water content, and pressure and temperature conditions of mylonitic 101
overprinting by Menegon et al. (2017), who concentrated on the deformation mechanisms that 102
allow localisation of viscous creep within pseudotachylyte in the lower crust. This current study 103
utilises these existing constraints but focuses on microstructural evidence for additional transient 104
non-steady state episodes of viscous deformation that appear to have occurred locally and 105
episodically within the ongoing mylonitisation characterised by Menegon et al. (2017). We constrain 106
the flow stress, strain rate and viscosity of these transient perturbations and discuss the ability of 107
pseudotachylyte to accommodate short-term variability in the conditions of solid-state viscous 108
creep. We attempt to relate deformation conditions recorded by microstructures in exhumed fault 109
rocks to creep rates and viscosities inferred from geodetic measurements during the postseismic and 110
interseismic stages of the seismic cycle, thus bridging the gap between the geological record and the 111
direct observation of earthquake cycle deformation.
112 113
2. Geological setting and sample description 114
The Lofoten Islands in northern Norway (Fig. 1a) are predominantly composed of Archaean to 115
Palaeoproterozoic ortho- and para-gneisses, intruded at 1.9-1.7 Ga by a suite of Anorthosite- 116
Mangerite-Charnockite-Granite (AMCG) bodies (Corfu, 2004). These intrusions are anhydrous, and 117
emplaced under granulite facies conditions, which were estimated at 750-800 °C and 0.4-1.2 GPa 118
(Markl et al., 1998). The primary igneous textures and the dry granulite facies mineral assemblages 119
are generally well preserved. Lofoten represents a tectonic window of basement rocks from the 120
Baltica plate that are usually buried beneath the Caledonian nappe pile. The Lofoten basement rocks 121
have largely escaped the Caledonian tectono-metamorphic overprint as the result of the limited 122
availability of fluids necessary to facilitate viscous deformation of the anhydrous, strong granulites 123
(Steltenpohl et al., 2004).
124
Granulite, eclogite, and amphibolite facies localised shear zones are common in Lofoten and are 125
frequently associated with large volumes of pseudotachylytes (Fig. 1b). Available evidence indicates 126
that these pseudotachylytes developed under lower crustal conditions and provided the necessary 127
weak precursors on which mylonitic shear zones subsequently localized (Menegon et al., 2017). The 128
Nusfjord anorthosite (Fig. 1a) is cut by a network of localised shear zones that contain both pristine 129
(Fig. 1b) and mylonitised (Fig. 1c-d) pseudotachylytes. Between these shear zones, the anorthosite is 130
undeformed, preserving a coarse-grained igneous texture with up to 1-5 cm large crystals of 131
plagioclase, associated with minor amounts (< 10 vol. %) of clinopyroxene ± amphibole ± quartz ± 132
orthopyroxene ± garnet ± biotite. Non-mylonitised, pristine pseudotachylytes (Fig. 1b) range in 133
thickness from < 1 cm to ca. 10 cm and can be followed along strike for up to 100 metres.
134
Mylonitised pseudotachylytes are recognised in the field by the local preservation of 135
pseudotachylyte breccia pockets and injection veins both along and across the mylonitic foliation 136
(Fig. 1c,d). They form shear zones that typically range in thickness from < 1 cm to ~30 cm, and that 137
are continuous along strike for over 1km on the eastern Nusfjord ridge. The thickest shear zones 138
may have derived from the cyclical mylonitic overprint of several generations of pseudotachylytes, 139
as indicated, at least locally, by the occurrence of pseudotachylyte veins discordant to the main 140
mylonitic foliation and only partially transposed into it (Fig. 1d). The shear zones also incorporate 141
parts of the damage zone around pseudotachylyte veins, or alternatively anorthosite, charnockite 142
and granite dykes along which pseudotachylyte veins have exploited the boundaries (Fig. 2a). This 143
results in shear zones that are somewhat wider than any pseudotachylyte veins they incorporate 144
(Fig. 2), but nevertheless are of limited width restricted by the precursor structure. Mylonitised and 145
non-mylonitised pseudotachylytes are observed adjacent to each other (Fig. 2b-d) in structural 146
contexts where faults decorated with non-mylonitised pseudotachylyte cut undeformed anorthosite 147
blocks between closely spaced (<20 m) mylonitised pseudotachylyte-bearing shear zones. These 148
non-mylonitised pseudotachylyte faults appear to represent frictional deformation of the strong 149
anorthosite host rock related to ongoing viscous creep on the shear zones and do not form in the 150
same orientation as the shear zones (Figs. 2c, d). Detailed field descriptions of the pseudotachylytes 151
and shear zones can be found in Menegon et al. (2017).
152
The shear zone network in the Nusfjord east region consists of three main sets of shear zones, all 153
containing variable amounts of mylonitised pseudotachylytes, although mylonitised 154
pseudotachylytes are most extensively developed in the ‘set 1’ shear zone orientation, dipping 155
steeply towards the SE with normal-oblique kinematics (Figure 2b, Menegon et al., 2017). Mutual 156
crosscutting relationships indicate that these three shear zone sets were broadly coeval. P-T 157
conditions of pseudotachylyte formation and mylonitisation were estimated at 650-750°C and 0.7- 158
0.8 GPa, corresponding to an approximate depth range of 24-30 km (Menegon et al., 2017). 40Ar-39 Ar 159
dating of amphibole from localised amphibolite facies shear zones in the Nusfjord area with similar 160
orientation to the Nusfjord east network yielded an age range of 433-413 Ma (Fournier et al., 2016;
161
Steltenpohl et al., 2003). Given the similarity of structures, mineral assemblages and metamorphic 162
conditions, this age range is also taken as representative of the Nusfjord east shear zone network 163
described here. The regional tectonic context at this time suggests that the Nusfjord shear zones 164
represent the rheological response of the Baltica basement underthrusting Laurentia during the 165
collisional (Scandian) stage of the Caledonian Orogeny.
166
Much of the primary structure of the mylonitised pseudotachylytes has been overprinted, so that it 167
is difficult to constrain any information about the nature of the seismicity that produced them. In 168
addition, any brittle fault system that could potentially have extended these shear zones to 169
structurally higher crustal levels is not preserved in the field around the Nusfjord area.
170
Pristine pseudotachylytes preserve a crystalline matrix consisting of plagioclase with some 171
combination typically of amphibole, clinopyroxene, orthopyroxene, biotite, and K-feldspar.
172
Pseudotachylyte veins may be enriched in biotite, K-feldspar and quartz, either in the crystalline 173
matrix or as unmelted clasts, where the pseudotachylytes cut through granitic and charnockitic pods 174
or dykes. Mylonitised pseudotachylyte - bearing shear zones contain sheared clasts of dismembered 175
and partially recrystallized plagioclase, sometimes in combination with partially transposed and 176
flattened injection veins (Fig. 3), and melt quenching related microstructures such as spherulites, 177
dendritic crystals and microlites (e.g. Figs. 3a-b) are not preserved. The amount of strain recorded in 178
the mylonitised pseudotachylytes varies from low, as indicated by the moderate inclination of 179
elongated survivor clasts with respect to the shear zone boundary (Fig. 2d in Menegon et al., 2017), 180
to very high, as indicated by the development of a strongly banded mylonitic foliation where clasts 181
of host rock are flattened and elongated parallel to the shear zone boundary (Fig. 2b). The 182
deformation microstructures analysed in this study have been observed in a large number of 183
samples of pseudotachylyte - bearing shear zones from Nusfjord east. Here we present the detailed 184
microstructural analysis of representative samples from shear zone ‘269’ (samples 269C and 269E2:
185
Figs. 3,4). The shear zone has the typical set 1 orientation and kinematics (dip and dip direction:
186
76/138) and includes a mylonitised pseudotachylyte vein exploiting the boundary of a charnockite 187
dyke (Fig. 2a), which is also variably sheared along strike. The shear zone can be traced 188
discontinuously for around ten metres and the samples were located approximately the same 189
distance along strike from each other.
190
3. Methods 191
3.1 Electron backscatter diffraction (EBSD) 192
The EBSD data were acquired using a JEOL 7001 FEG-SEM equipped with a NordLysMax detector at 193
the Electron Microscopy Centre of the University of Plymouth. Polished thin sections were carbon 194
coated prior to EBSD analysis. Working conditions during acquisition of the EBSD patterns were 20 195
kV accelerating voltage, 12 nA probe current, 70° sample tilt, ~ 20 mm working distance and high 196
vacuum. EBSD patterns of quartz were acquired and indexed with the AZtec software (Oxford 197
Instruments) on four rectangular grids from two samples of mylonitised pseudotachylytes, 269C and 198
269E2 (Fig. 3c,e, Fig. 4), with step sizes ranging from 0.4 µm to 1.2 µm and using the “quartz-new”
199
match unit from the HKL-Oxford Instruments’ database. EBSD data were processed using the 200
Channel 5 software (Oxford Instruments).
201
EBSD data are presented as Grain Orientation Spread (GOS) maps, as quartz c-axis pole figures, and 202
as grain size distribution histograms. The GOS is a measure of the internal strain of a grain defined as 203
the average misorientation angle between each pixel in a grain and that grain’s mean orientation 204
(Wright et al., 2011). The orientation of quartz c-axis was plotted on contoured pole figures (lower 205
hemisphere of the equal angle projection) as one point per grain. The pole figures are oriented in 206
the plane containing the pole to the mylonitic foliation (the Z-axis) and the stretching lineation (X- 207
axis). The orientation of the X-axis and Z-axis is shown on each GOS map. The grain size distribution 208
histograms were plotted as frequency (in %) vs grain size (in m) calculated as the diameter of the 209
circle with an area equivalent to that of the grain.
210
3.2 Flow stress, strain rate and viscosity calculations 211
The EBSD-calibrated recrystallized grain size piezometer for quartz of Cross et al. (2017) was used to 212
calculate the flow stress of mylonitisation. Recrystallized quartz grains were distinguished from the 213
relict grains by their Grain Orientation Spread (GOS). The critical GOS value corresponding to the 214
threshold between relic grains and recrystallized grains was established from the ‘kneepoint’ in the 215
cumulative frequency plot of GOS for the entire quartz population (Fig. S1). The root mean squared 216
grain size (DRMS) of the recrystallized quartz grains was used in the ‘sliding’ piezometer of Cross et al.
217
(2017) which states D = 104.22±0.51 σ-1.59±0.26, where D is the mean grain size and σ is the flow stress.
218
The dislocation creep flow laws for quartz of Hirth et al. (2001) and Tokle et al. (2019) were used to 219
calculate and compare possible strain rates for quartz, based on the deformation microstructures 220
characterised in section 4.1. For the recrystallized pseudotachylyte matrix, the wet diffusion creep 221
flow law for an anorthite-diopside aggregate (Dimanov & Dresen, 2005) was chosen as appropriate 222
due to the fine grain size and lack of crystal preferred orientation (CPO) in Nusfjord 223
pseudotachylytes, as suggested by Menegon et al. (2017). In Nusfjord, the pseudotachylytes are 224
relatively hydrated relative to the host rock and the H2O content of the recrystallized 225
pseudotachylyte is 0.2-0.4 wt. % (Menegon et al., 2017), above that of the ‘wet’ samples (0.05 wt%) 226
of Dimanov and Dresen (2005) and hence the wet flow law was chosen. The H2O content of the 227
experiments compiled in Tokle et al. (2019) sit in the range 0.1-0.4 wt % and are therefore very 228
similar to these Lofoten samples.
229
To represent deformation of the entire pseudotachylyte - bearing shear zone, a simple mixed flow 230
law was calculated for 10% quartz deforming by dislocation creep and 90% anorthite-diopside 231
deforming by diffusion creep, in the form 𝜀̇𝑚𝑖𝑥𝑒𝑑= [𝑋𝑞𝑢𝑎𝑟𝑡𝑧𝜀̇𝑞𝑢𝑎𝑟𝑡𝑧+ 𝑋𝐴𝑛−𝐷𝑖𝜀̇𝐴𝑛−𝐷𝑖] where 232
𝜀̇𝑚𝑖𝑥𝑒𝑑 is the strain rate of the mixed flow law, X denotes the fraction and 𝜀̇ the strain rate of quartz 233
and anorthite-diopside respectively. The strain rates were multiplied by √3 to convert from axial to 234
shear strain rates (Paterson and Olgaard, 2000). 10% is perhaps an overestimate of quartz content in 235
pseudotachylyte - bearing shear zones in Nusfjord east, but the mixed strain rate proves insensitive 236
to small adjustments in quartz volume.
237
Effective viscosity of the pseudotachylyte-bearing shear zones was calculated from the flow stress 238
and the strain rate (using the Hirth et al., 2001 flow law to model the contribution of dislocation 239
creep of quartz) as η = σ/2𝜀̇, where σ is the flow stress and 𝜀̇ the strain rate. Errors were propagated 240
through the process from the uncertainties in temperature and pressure (as stated in Menegon et 241
al., 2017) as well as the errors published for parameters in the flow laws and piezometric relations.
242 243
4. Results 244
4.1 Microstructure, crystallographic preferred orientation, and grain size of quartz in 245
pseudotachylyte - bearing shear zones 246
Monomineralic quartz domains are locally present within shear zones in different textural positions:
247
(1) in deformed host rock domains at the immediate margins to mylonitised pseudotachylyte veins 248
in mylonitised breccias (Figs. 1c, 3c-d, 4a-f), (2) in ribbons entirely wrapped by the recrystallized 249
pseudotachylyte matrix within mylonitised brecciated networks, (Figs. 3e-f, 4g-h), and (3) as thin (50- 250
200 m thick), elongate ribbons within sharp mylonitised veins derived from the pseudotachylyte 251
generation surfaces (Fig. S2a). In all these domains, quartz occurs as recrystallized fine grains 252
mantling larger grains. The large grains show undulose extinction and shape elongation (Fig. 4, b,d, 253
Fig. S3), and contain a high density of low-angle boundaries that define subgrains of the same size of 254
the recrystallized grains (Fig. 5). The smaller grains have low Grain Orientation Spread (GOS) values 255
(typically ≤4.3°, Fig. 5), which are used to separate the populations of recrystallized grains from the 256
larger relict grain population (Cross et al., 2017). Where quartz occurs further away (Figs. 3g-h) from 257
the sheared pseudotachylyte veins, it lacks both the shape elongation and the frequency of finer 258
grains that is observed in the immediate vicinity to the veins. Quartz, plagioclase and clinopyroxene 259
in the proximal damage zone flanking pristine pseudotachylytes do not show evidence of crystal 260
plastic deformation and dynamic recrystallization (Figs. 3a, b).
261
The <c> axis pole figures for the recrystallized and relict quartz grains are very similar at each 262
analysed site (Fig. 5). In sample 269C, <c> axis orientations form an incomplete single girdle with a 263
maximum near the centre of the pole figure. In sample 269E2, at analysis sites A and C, the <c> axes 264
form maxima oblique to the mylonitic foliation (Fig. 5).
265
In sample 269E2, site B (Fig. 6) shows a region characterised by well-developed quartz triple 266
junctions, equant grain shape, low GOS and a larger grain size than the rest of the recrystallized 267
population, similar to foam textures described in Kidder et al. (2016). These grains have a similar 268
crystal preferred orientation (CPO) to the relict and finer recrystallized quartz aggregates preserved 269
locally elsewhere around the margin of this vein (Figs. 5,6). Similar features are seen very locally 270
within site C (Fig. 5).
271
Within mylonitised pseudotachylyte veins, quartz locally occurs also as dispersed grains within the 272
vein matrix (Figs. 5 - 7). This appears to be the result of the progressive disaggregation of 273
polycrystalline recrystallized ribbons leading to phase mixing (Cross & Skemer, 2017, Kilian et al., 274
2011). One-grain-wide elongate ribbons parallel to the vein margins are progresisvely detached 275
from the margin quartz aggregates (Fig. 7), and individual quartz grains are separated from the main 276
ribbon by grains of other phases (mostly plagioclase). The crystallographic preferred orientation of 277
the disseminated quartz grains is very weak or absent (Fig. 7). The foam-texture regions, the 278
disseminated quartz in the vein, and the thinnest recrystallized ribbons (only 1-2 grains in width) 279
were excluded from grain size calculations used for piezometry due to the effects of annealing and 280
second phase pinning on the grain size (Herwegh et al., 2011).
281
In summary, we interpret the quartz microstructure as indicative of dislocation creep and dynamic 282
recrystallization by dominant subgrain rotation, based on the occurrence of core-and-mantle 283
microstructures, on the similar size of subgrains and of recrystallized grains, and on the overlap in 284
the c-axis crystallographic preferred orientation between relict grains and recrystallized grains (Fig.
285
5). The similarity between subgrain size and recrystallized grain size indicates that grain growth after 286
subgrain rotation recrystallization was negligible. On the other hand, the local development of foam 287
texture quartz is likely to be the result of static annealing.
288
The root-mean-squared (RMS) grain sizes of the recrystallized quartz populations are 12.5 μm, 12.4 289
μm, and 9.2 μm for samples 269C, 269E2 site A and 269E2 site C respectively. Using the sliding 290
piezometer calibrated by Cross et al. (2017), these recrystallized quartz populations suggest flow 291
stresses of 92.0 MPa, 92.3 MPa, and 111.7 MPa, giving a mean flow stress across these samples of 292
98 MPa ± 18 MPa (2σ).
293
4.2 Estimate of strain rates in pseudotachylyte - bearing shear zones 294
The rheology of quartz-bearing pseudotachylyte shear zones in Nusfjord east was modelled using a 295
mixed flow law that considers a 10% volume fraction of quartz deforming by dislocation creep (flow 296
laws from Hirth et al., 2001 and Tokle et al., 2019) and a 90% volume fraction of anorthite-diopside 297
aggregate deforming by wet diffusion creep (flow law from Dimanov & Dresen, 2005), based on the 298
typical proportions of quartz in and around the mylonitised pseudotachylytes (see section 3.2).
299
At temperatures of 700°C and flow stress of 98 MPa, the predicted strain rates for all flow laws fall 300
within the range of 10-10 – 10-7 s-1 (Fig. 8). At this bulk flow stress, wet dislocation creep of quartz 301
allows deformation at a faster strain rate than wet diffusion creep of anorthite-diopside.
302
Mylonitisation of the pseudotachylyte involved concurrent dislocation creep of the quartz ribbons 303
and grain size sensitive creep of the anorthite-diopside matrix, similar to other deformed and 304
recrystallized pseudotachylytes (Price et al., 2012; White, 1996). We use the flow stress derived from 305
quartz piezometry to calculate all strain rates, including in the mixed flow law which combines the 306
behaviour of quartz and the pseudotachylyte matrix (the evidence behind this assumption of 307
constant stress is discussed in section 5.1). At a flow stress of 98 MPa, the mixed flow law using the 308
Hirth et al., 2001 quartz dislocation creep flow law results in a strain rate of 6.3 x 10-9 s-1. Quartz 309
dislocation creep flow laws are also shown from Tokle et al. (2019), with stress exponents of n=3 and 310
n=4; using these to represent quartz in a mixed flow law at a stress of 98 MPa results in strain rates 311
of 1.4 x 10-9 s-1 and 3.4 x 10-9 s-1 respectively. Since these resultant mixed flow law strain rates are 312
very similar, only the mixed flow law using the Hirth et al. (2001) quartz dislocation creep flow law is 313
used in subsequent calculations of viscosity. The choice of these flow laws is discussed in section 5.1.
314
The mixed flow laws, being combined from both diffusion creep (linear viscous) and dislocation 315
creep (power law) flow laws, are inherently non-linear.
316
4.3 Estimate of the effective viscosity of mylonitised pseudotachylyte 317
The effective viscosity of the Nusfjord mylonitised pseudotachylytes deforming at 700°C is 3.6 x 1016 318
Pa.s using the strain rate of 6.3 x 10-9 s-1 derived from themixed quartz + anorthite-diopside flow 319
law, and the flow stress of 98 MPa. When compared to equivalent temperature estimates from 320
geodetic and geological studies of the viscosity of the lower crust (Fig. 9), this result is comparable 321
with the lowest values reported, although most literature values are at least an order of magnitude 322
higher. The lack of viscous deformation observed in the anorthosite outside the shear zones suggests 323
that much higher stresses than ever were present would be needed to induce viscous creep in the 324
anorthosite. Hence, it would be unrealistic to combine an anorthosite viscosity with the mylonitised 325
pseudotachylyte viscosity to represent the entire ~1 km crustal width encompassed by the shear 326
zone network. Additionally, the anorthosite viscosity under dry dislocation creep (Dimanov & Dresen 327
2007) would be several orders of magnitude higher than reported literature values, being around 328
1024 Pa.s (Fig. 9). In order to ensure that they represent deformation of fault zones at depth, all 329
geodetic values shown in Fig. 9 are derived from postseismic observations. The geological studies do 330
not necessarily identify any association with seismogenic indicators, but are taken from observations 331
of shear zones and hence are inherently localised.
332
5. Discussion 333
5.1 High strain rate transients preserved in the quartz recrystallized grain size 334
Recrystallized quartz in the mylonitised pseudotachylytes (Figs. 4, 5, S2, S3) presents microstructures 335
typical of bulging and subgrain rotation recrystallization. These microstructures are more generally 336
associated with deformation temperatures up to 500°C (e.g. Stipp et al., 2002) rather than the 700 ± 337
50° C established for the viscous deformation of the Nusfjord east pseudotachylyte (Menegon et al., 338
2017). However, there is no evidence that quartz recrystallization occurred during a lower 339
temperature overprint, in that recrystallized quartz coexists with neoblasts of clinopyroxene (Fig.
340
S2b), pargasitic amphibole and calcic plagioclase in the mylonitic foliation (Menegon et al., 2017).
341
Additionally, the c-axis orientations of the fine- and coarse- grains are the same (Figs. 5,6), whereas 342
if the recrystallized grains were re-deformed by some later event, some differences in the CPO might 343
be expected. Any undetected lower temperature viscous overprint of the pseudotachylytes seems 344
unlikely, as the temperature estimate was derived from thermodynamic modelling of the whole 345
pseudotachylyte vein assemblage and from amphibole-plagioclase geothermobarometry using 346
amphibole grains that grew into synkinematic dilatant sites (Menegon et al., 2017). Fine 347
recrystallized grain size of quartz in lower crustal shear zones has been attributed to dry conditions 348
during deformation that inhibit efficient recovery (Fitz Gerald et al., 2006; Menegon et al., 2011), 349
which would be one explanation for the occurrence of ‘low-temperature’ quartz microstructures at 350
700°C. However, the Nusfjord mylonitised pseudotachylytes deformed under locally H2O-present 351
conditions which facilitated the nucleation of amphibole along grain- and phase boundaries 352
(Menegon et al., 2017). Thus, we interpret the apparent ‘low temperature’ microstructure of the 353
fine-grained recrystallized quartz to instead be the result of high stress and strain rate deformation 354
at the same P-T conditions of 700°C and 0.8 GPa suggested by Menegon et al. (2017) for the long- 355
term viscous creep in the mylonites. This reinforces that care should be taken when qualitatively 356
attributing deformation temperatures to microstructures due to the accompanying trade-off 357
between strain rate, water content and stress (Hirth & Tullis, 1992, Piazolo et al., 2002, Tokle et al., 358
2019).
359
Two published flow laws are compared in Fig. 8 for dislocation creep of quartz and the resultant 360
mixed flow laws (Hirth et al., 2001, Tokle et al., 2019). Both have been chosen for their attempts to 361
reconcile experimental with natural deformation conditions and inclusion of microstructural 362
comparisons. The more recent work of Tokle et al. (2019) suggests that the stress exponent of quartz 363
dislocation creep flow laws may change with changing deformation temperatures and stresses – at 364
high stress and low temperatures, n = 3 is deemed most applicable, and at low stress and high 365
temperatures, n = 4. Around the calculated differential stress of 98 MPa, all the quartz flow laws in 366
fact give a mixed flow law strain rate on the order of 10-9 s-1 (Fig. 8), so we are confident that this 367
figure is not biased by the choice of published flow law. Under these deformation conditions, the 368
Tokle et al. (2019) flow law using a stress exponent (n) of 4 predicts that prism <a> will be the rate- 369
limiting slip system, whereas basal <a> is predicted to be the rate-limiting slip system for the n = 3 370
flow law. In the Nusfjord samples, 269C shows a CPO maxima near to the Y-axis (Fig. 5), consistent 371
with prism <a> slip. However, the CPOs for the analysed sites in sample 269E2 are less clear but 372
more suggestive of a mix of basal <a> and rhomb <a> slip. At 98 MPa flow stress, as calculated from 373
the quartz piezometry, the n = 4 flow law is marginally faster (weaker) than the n = 3 flow law (Tokle 374
et al., 2019), so the n = 4 slip on prism <a> would be expected to occur if the conditions made that 375
feasible. Tokle et al. (2019) also state that the n = 4 flow law is applicable under low stress (and/or 376
high temperature) conditions, and the n = 3 law under higher stress (and/or lower temperature). It is 377
difficult to extrapolate the experimental temperatures and pressures to compare with our natural 378
conditions, but we suggest that the difference in CPO between samples 269C and 269E2 do not 379
represent a stress difference, based on the similar piezometry results, and so do not warrant the 380
attribution of separate flow laws with different stress exponents. With reference to the choice of 381
flow law to represent the pseudotachylyte matrix, although amphibole is also present, its CPO 382
suggests that it deformed via diffusion creep (Menegon et al., 2017). Amphibole deforming by 383
diffusion creep under lower crustal temperature and pressure conditions displays similar strength to 384
wet anorthite deforming by diffusion creep (Getsinger & Hirth, 2014). Hence, we prefer to use the 385
established anorthite-diopside flow laws (Dimanov & Dresen, 2005) rather than attempt to account 386
for amphibole behaviour.
387
Calculation of strain rates for the anorthite-diopside pseudotachylyte matrix, and of the combined 388
flow law, assumes that the flow stress calculated from quartz piezometry is homogeneous across the 389
vein. Although the deformation mechanisms in the recrystallized polymineralic pseudotachylyte 390
matrix and in the monomineralic quartz aggregates were different (diffusion creep and grain 391
boundary sliding in the recrystallized pseudotachylytes vs. dislocation creep in quartz), there does 392
not seem to be any microstructural evidence for a significant viscosity contrast between the quartz 393
where it occurs in monomineralic ribbons within the vein and the matrix of the pseudotachylyte 394
(Figs. 3e-f, 4a,e,g), whereas pinch-and-swell, buckling and boudinage microstructures might be 395
expected if there were. These microstructures are expected to form even at small strains (Gardner 396
et al., 2016). In the absence of clear evidence of viscosity contrasts between monomineralic 397
recrystallized domains and polymineralic matrix, recrystallized grain size palaeopiezometry has 398
yielded representative results in the study of shear zone rheology elsewhere in the lower crust (e.g.
399
Mehl & Hirth, 2008; Viegas et al., 2016; Wex et al., 2019), and here we adopt the same approach. In 400
addition, in the context of Voigt-Reuss elastic limits, this constant-stress state indicates the lower 401
(Reuss) bound for the bulk viscous strength of the pseudotachylyte vein (Hill, 1965), but noting the 402
similarity in strain rates for quartz and for anorthite-diopside aggregates at the predicted flow stress 403
(Fig. 8), the upper Voigt bound (constant strain rate) is not expected to be significantly different.
404
Hence, either both strain rate and stress were approximately constant across the vein, or both must 405
vary in complementary ways to maintain the constant bulk viscosity.
406
We consider the quartz deformation outlined here as representative of the transient high stress 407
deformation localised in the pseudotachylyte - bearing shear zones. This is supported by the 408
observation that the systematic occurrence of fine-grained recrystallized populations is limited to 409
the proximity to the mylonitised pseudotachylyte veins or to their interior (compare Figs. 3g-h and 410
Fig. 4), arguing for highly localised enhanced strain rate in the mylonitised pseudotachylytes. In 411
addition, the quartz deformation cannot pre-date the pseudotachylyte formation, because the 412
quartz in the damage zone of unmylonitised pseudotachylytes is also undeformed (Fig. 3b). Although 413
quartz is present only in limited amount (ca. 10 vol. %) within and around the analysed mylonitised 414
pseudotachylyte veins, the low variation of recrystallized grain sizes and resultant flow stresses for 415
samples from different shear zone strands suggests that discontinuous aggregates of quartz are able 416
to consistently capture the high strain rate deformation localised to the sheared pseudotachylytes.
417
Support for the short-term transience of the high stress, high strain rate state is provided by local 418
low stress overprinting of the fine-grained recrystallized quartz. In sample 269E2 site B (Fig. 6), 419
quartz at the immediate margin of the mylonitised pseudotachylyte forms a band of much coarser 420
grains than the adjacent recrystallized quartz, and has lower internal misorientation than the larger 421
relict grains contained in the same region. Equant grains and 120° triple junctions are also common 422
in this coarser quartz band (Fig. 6). The quartz in this immediate margin region have a CPO with 423
similar orientation of <c> axes to the relict and fine recrystallized quartz populations of the more 424
distal margin, forming maxima moderately oblique (in a clockwise direction) to the trace of the 425
foliation with some dispersion extending out towards the Y axis. This preservation of the CPO during 426
annealing, combined with a slight weakening of the CPO, is consistent with the findings of quartz 427
static annealing studies (Heilbronner & Tullis, 2002). A partial development of a similar 428
microstructure is also observed in the margin to the pseudotachylyte vein in sample 269E2 site C 429
(Fig. 5), where some of the recrystallized grains with low GOS values are larger than the typical 430
subgrain size in this sample, and also display 120° triple junctions. These features are classified as a 431
foam-texture, which has been shown to develop in quartz as a response to a rapid decrease in stress 432
(Kidder et al., 2016). The grains at the edges of the foam –texture domains appear to be growing 433
over the domains of relict and fine-grained recrystallized quartz (Figs. 5,6). We interpret this contrast 434
between the finer-grained and the foam-texture recrystallized quartz populations to represent two 435
stages of progressive deformation, as indicated by the consistent CPO. Partial overprinting of the 436
fine-grained recrystallized population by the foam-texture quartz may represent a localised record of 437
the stress decrease during the creep following the high stress transient (e.g. Trepmann et al., 2007).
438
The high stresses and strain rates recorded from the finer recrystallized quartz are therefore a 439
snapshot of temporary transient deformation conditions, in line with recent experimental results 440
which demonstrated the ability of quartz microstructures and grain size to capture transients (Kidder 441
et al., 2016). Our interpretation, therefore, is that in the Nusfjord mylonitised pseudotachylytes, the 442
relict and fine-grained recrystallized quartz populations record an initial transient high stress, high 443
strain rate deformation event localised within the sheared pseudotachylyte.
444
The flow stress of 98 MPa recorded in the recrystallized grain size of fine-grained quartz in 445
mylonitised pseudotachylytes is higher than the differential stresses typically estimated for shear 446
zones at lower crustal depths (Behr and Platt, 2014; Getsinger et al., 2013), but transient high 447
stresses hundreds of megapascals higher than steady-state have been observed and modelled for 448
natural shear zones near the frictional-viscous transition (Ellis & Stöckhert, 2004; Kuster & Stöckhert, 449
1999; Trepmann & Stöckhert, 2003; Matysiak & Trepmann, 2012). This scenario fits with the likely 450
depth range and nature of viscous deformation on the Nusfjord east shear zones which were likely 451
to have been active at depths close to the frictional-viscous transition for dry plagioclase (Fig. 10a), 452
typically occuring at depths > 20km. At such depths, brittle seismic failure of the dry anorthosite host 453
rock requires differential stresses in excess of hundreds of megapascals (Fig. 10a). Hence, even when 454
the stress drop is nearly complete (i.e. for / in the range of 0.6-0.9, where is the earthquake 455
stress drop and is the maximum shear stress on the seismogenic fault plane in the pre-earthquake 456
stress field; Hardebeck and Okada, 2018), residual differential stresses during the postseismic period 457
could be expected to be on the order of 100 MPa after a ~ 1 GPa coseismic stress drop. Whilst long- 458
term, steady state viscous creep localised on the shear zones (including along the mylonitised 459
pseudotachylytes) in Nusfjord may generally have taken place at more typical geological strain rates 460
of 10-15-10-13 s-1 (Fagereng & Biggs 2018) under relatively low differential stresses during the 461
interseismic period (i.e. blue curves in Fig. 10a), transient higher differential stresses and strain 462
rates, such as 10-9 s-1 derived in this study, can also be supported by localised viscous creep in the 463
same material (orange curve in Fig. 10a).
464
It has been suggested that transient high strain rate deformation and steady state low strain rate 465
deformation may need to be accommodated in mechanically different materials, as proposed in 466
relation to the postseismic behaviour of the 1992 Landers earthquake (Ivins, 1996). However, given 467
the high strength contrast between the pseudotachylyte shear zones and the surrounding, largely 468
undeformed anorthosite, it is difficult to envisage long-term steady-state creep not being localised 469
onto pseudotachylyte - bearing shear zones in Nusfjord. Pseudotachylyte - bearing shear zones were 470
therefore able to support both low viscosity creep under transient high strain rates and higher 471
viscosity deformation during steady state deformation, as indicated by localised overprint of lower 472
stress foam-textured quartz. It may be that the quartz microstructures in the Nusfjord 473
pseudotachylytes were predominantly preserved in the high stress state and recorded the lower 474
stress stage only locally. The general preservation of high stress quartz microstructures is in 475
accordance with experiments that suggest that quartz deformed at high stress and strain rate 476
exhibits strong strain hardening (Hobbs, 1968), making it less likely to progressively change 477
microstructure under subsequent reductions in stress or strain rate.
478
5.2 Timing of transient high stress, high strain rate deformation in the Nusfjord pseudotachylytes 479
Although the rheology of pseudotachylyte - bearing shear zones in the lower crust has been shown 480
here to support short term episodes of elevated stress and strain rates during non-linear viscous 481
creep, there are two scenarios which could be proposed for when these transient conditions 482
occurred. Firstly, the high stress and strain rate transients could be part of the same seismic cycle 483
which generated the pseudotachylytes, i.e. occurs almost immediately on crystallisation of the 484
pseudotachylyte on the cessation of coseismic slip. Alternatively, the pseudotachylytes could be 485
somewhat or even very much older than a deformation event that generated high stresses and 486
strain rates during localised viscous creep in the lower crust. Such an event could have potentially 487
been the seismogenic reactivation of an overlying upper crustal fault zone, with a downward 488
propagation of high stress in the lower crust (Ellis and Stöckhert, 2004). The timing of transient high 489
strain rate, high differential stress creep cannot be easily constrained in the Nusfjord 490
pseudotachylytes, but considering a cooling model and likely crystallization rates for these 491
pseudotachylytes (Fig. S5) suggests that solid-state viscous deformation could be accommodated 492
within minutes to a couple of hours after the seismicity which generated these veins (e.g. Ferrand et 493
al., 2018). Given that pseudotachylytes in the Nusfjord shear zones were generated episodically 494
within the period of long-term background viscous creep (Fig. 1d), we propose that the observed 495
transient high stress, high strain rate state (indicated by the fine-grained quartz) and subsequent 496
lower stress overprint (indicated by the foam texture quartz) represent variations over a timescale of 497
an earthquake cycles (Fig. 10b). The high differential stress of 98 MPa is not envisaged to be the 498
peak stress experienced during the seismic cycle, because the solid-state viscous deformation 499
recorded in the fine-grained recrystallized quartz must take place only after crystallization of the 500
pseudotachylyte melt has occurred and so the highest, immediately post-earthquake stress will have 501
already decayed somewhat; however, significant stress perturbations in the lower crust are perhaps 502
experienced for ~ 100 years or more (Ellis & Stöckhert, 2004). The strain rate of 10-9 s-1, as recorded 503
in these samples, is rare (or rarely preserved) in geological studies, with ‘typical’ strain rates being 504
10-13-10-15 s-1 (Fagereng & Biggs, 2018). On the geologically short timescales of earthquake cycles, the 505
most likely driver of rapid aseismic strain rates is postseismic relaxation (Fig. 10b).
506
5.3 Postseismic relaxation as a potential driver of transient high stress and strain rate 507
Postseismic relaxation characterises the ~100 year period immediately following seismic slip where 508
rapidly relaxing high strain rates and stresses, and low but increasing crustal viscosities, are observed 509
on and around the fault zone until some long-term interseismic steady-state is reached (Thatcher, 510
1983, Ingleby & Wright, 2017). Hence, if some of this postseismic relaxation is accommodated via 511
viscous creep, any resultant deformation microstructure and recrystallisation is expected to 512
progressively evolve into a lower stress, lower strain rate deformation phase, potentially 513
overprinting the transient high strain rate microstructures (Kidder et al., 2016; Trepmann &
514
Stöckhert, 2003). In the Nusfjord psedutoachylytes we see this expressed in the local development 515
of foam-texture quartz. In order for the foam texture quartz (Fig. 6) to feasibly represent an 516
interseismic component of the seismic cycle (Fig. 10b), the grain growth and annealing of quartz 517
must also be able to produce a maximum grain size of ~ 30 µm from a minimum initial grain size of 518
~6.5 µm (Fig. 6) within a timescale applicable to seismic cycles. Parameters for the static grain 519
growth of quartz are available from Wightman et al. (2006), based on natural quartz samples, and 520
from Fukuda et al. (2019), based on quartz grown under experimental conditions. Grain growth is 521
modelled on the concept dn – d0n = kt, where d is the grain size at time t, d0 the initial grain size, k the 522
rate constant and n is the growth exponent (e.g. Fukuda et al., 2019). Using temperatures of 650- 523
750°C and pressures of 0.7 – 0.75 GPa for steady-state deformation in the Nusfjord mylonitised 524
pseudotachylytes (Menegon et al., 2017), and water fugacities calculated for these conditions 525
(Tödheide, 1972), the minimum period necessary for transformation from the fine-grained 526
recrystallized quartz to the foam texture ranges from 1-100 years, inversely proportional to 527
temperature (Fig. 10c). The grain growth parameters of Wightman et al., 2006 are somewhat 528
pressure dependent but those of Fukuda et al. (2019) are not, excepting the input of water fugacity.
529
Although the more recent work of Fukuda et al. suggested that the temperature dependence of the 530
Wightman et al. (2006) grain growth parameters might be overestimated, the timescales estimated 531
by both methods are very similar under the Nusfjord deformation conditions. The 100 year timescale 532
given by both estimates is completely compatible with the periodicity of seismic cycles.
533
In addition, the non-linear mixed flow law that accounts for the combination of dislocation creep of 534
quartz plus diffusion creep of the pseudotachylyte matrix is one of the accepted rheologies that can 535
account for the observed rapid rates of postseismic strain rate decay (Ingleby & Wright, 2017). Such 536
non-linear creep may in fact be characteristic of pseudotachylytes where the conditions allow 537
viscous deformation via dislocation creep of the coarser grained survivor clasts and surrounding 538
damage zone in addition to the typical diffusion creep of the finer-grained pseudotachylyte matrix 539
(Passchier, 1982, Menegon et al., 2017). Although the pseudotachylytes would remain able to 540
viscously accommodate subsequent high stress and strain rate epsiodes for as long as they remained 541
fine-grained at similar P-T conditions, it seems unlikely that they would not also accommodate the 542
postseismic transients related to the coseismic event which formed them in the first place. In 543
conclusion, we suggest that the most likely origin of the observed stress and strain rate transients 544
was postseismic relaxation and that the pseudotachylytes locally record the changing stresses and 545
strain rates across a singe seismic cycle.
546
This study provides the first examination of recrystallized pseudotachylyte as an accommodator of 547
rapid postseismic creep. However, glassy pseudotachylyte is also proposed to undergo viscous flow 548
at temperatures found in the mid-crust and below, providing that the glass is wet (Proctor et al., 549
2017). Under these conditions, glassy pseudotachylyte experimentally shows low viscosities around 550
1012 Pa.s., several orders of magnitude lower than the viscosity predicted for recrystallized 551
pseudotachylytes in Lofoten. However, the experimental samples of Proctor et al. (2017) differ 552
widely from the Lofoten samples and many other exhumed natural pseudotachylytes, as they 553
contain <1% crystals, no clasts, have pore water, and the mechanism of deformation is not 554
intracrystalline plasticity and diffusion creep but flow of an amorphous matrix above the glass 555
transition temperature (Proctor et al., 2017). Whilst the conditions of deformation are different, the 556
temperature control on where pseudotachylytes can accommodate postseismic creep is highly 557
significant. At low temperatures equivalent to those in the upper crust, pseudotachylytes will be 558
brittle and often as strong as, or stronger than, the surrounding rock, and will not localise later 559
frictional failure after the coseismic event (Di Toro & Pennacchioni, 2005; Mitchell et al., 2016;
560
Proctor et al., 2017). Thus, the proposed accommodation of postseismic deformation by viscous 561
creep within pseudotachylytes is viable only in the mid- to lower- crust where thermally-activated 562
creep mechanisms are dominant.
563 564
5.4 Lower crustal viscosity and extent of localisation 565
Several geodetic observations of fault zone deformation at elevated strain rates are best explained 566
by some contribution from creep on localised lower crustal shear zones (Kenner & Segall, 2003;
567
Yamasaki et al., 2014), and both geological and numerical studies indicate that localisation is 568
generally expected at depth where the lower crust is relatively strong (Getsinger et al., 2013;
569
Montési, 2004). This scenario is consistent with what we infer for Nusfjord east, where the transient 570
high strain rate viscous creep was almost entirely localised on pseudotachylyte - bearing shear zones 571
(Figs. 2,3). The contribution to deformation from the wider coarse-grained anorthosite appears to 572
have been insignificant, with little evidence from field observations that any deformation has 573
occurred within it. Very little of the anorthosite can therefore have contributed to the deformation 574
and bulk strain rate across the kilometre-scale shear zone network. Instead, viscous deformation 575
must have continuously been highly localised onto the pseudotachylyte - bearing shear zones.
576
Although the viscosity of the pseudotachylyte (~1016 Pa.s) during the high strain rate deformation is 577
low relative to non-deforming or low strain rate lower crustal values, it is similar to some geodetic 578
postseismic lower crustal as well as some localised geological estimates (Fig. 9). The slight difference 579
in magnitude between geodetic postseismic viscosities and the value derived here may be due to the 580
difference in sampling scale between geodetic and microstructural studies, as well as to the high 581
level of localisation seen in the Nusfjord shear zones which may not be exactly comparable to some 582
mature continental scale faults sampled in geodetic studies. Generally, however, the similar 583
viscosities suggest that high strain rate deformation, such as is experienced during post-seismic 584
relaxation, is in fact to some extent localised where it is accommodated within the lower crust, 585
especially where the lower crust is known to be dry. This finding, alongside other recent work 586
(Hussain et al., 2018; Ingleby & Wright, 2017; Yamasaki et al., 2014), highlights that postseismic 587
studies of lower crustal viscosity - which are inherently undertaken in the vicinity of large fault zones 588
- do not derive estimates of the viscosity of the normal lower crust. They instead likely reflect the 589
heterogeneous viscosity across deep shear zones and the temporally varying viscosities induced by 590
the initial earthquake. In light of this, we suggest that although our viscosity estimates are derived 591
from narrow, highly localised shear zones, they nonetheless would be close to the viscosity that 592
would be seen across the deep roots of the Nusfjord east shear zone network during the active high 593
strain rate deformation by a geodetic survey, because of the extent of localisation and lack of 594
contribution to viscous creep from the host rock.
595
We highlight that, regardless of whether lower crustal pseudotachylytes are immediately 596
overprinted by viscous creep after their solidification from coseismic melts, or result from earlier 597
deformation events but are reactivated by later viscous creep, they represent weak domains that 598
could facilitate high strain rate deformation within the granulitic lower crust typical of thick 599
continental interiors. If viscous creep immediately follows pseudotachylyte generation, as is shown 600
to be possible from their rapid crystallization and cooling times, then additional effects such as the 601
introduction of water into the dry and strong lower crust (Jamtveit et al., 2018) will amplify the local 602
weakening effect.
603 604
6. Conclusions 605
Deformation recorded by pseudotachylyte - bearing mylonitic shear zones in Nusfjord east, Lofoten, 606
indicates relatively high differential stress and rapid strain rates. The strain rate in particular, at ~ 10- 607
9 s-1, is several magnitudes faster than typical geological processes. Such elevated strain rates are 608
interpreted to result from non steady-state flow associated with transient deformation of the lower 609
crust, presumably during postseismic relaxation. Viscosities of deforming pseudotachylyte - bearing 610
shear zones indicate similar values to transient lower crustal observations derived from geodetic 611
studies on active fault zones, supporting the inference that transient high strain rate creep can be 612
accommodated within lower crustal mylonitised pseudotachylytes. The strength contrast between 613
the fine-grained pseudotachylyte and the surrounding anorthosite causes localisation of the high 614
strain rate, high stress deformation, and this is likely to be the case in many lower crustal shear 615
zones hosted in dry, feldspar-rich lithologies.
616 617
Acknowledgements 618
This work was supported by the UK Natural Environment Research Council [grant number 619
NE/P001548/1 “The Geological Record of the Earthquake Cycle in the Lower Crust”]. We thank 620
Sandra Piazolo and Andrew Cross for their thorough and constructive reviews. The staff at the 621
Plymouth University Electron Microscopy Centre are thanked for support during SEM analysis. We 622
thank Tim Wright and Åke Fagereng for constructive discussion throughout the process of this study 623
and for their friendly reviews of the manuscript, as well as Jean-Philippe Avouac and Christie Rowe 624
for their constructive comments to an earlier version of the manuscript. According to the NERC data 625
management policy, data is available at the British Geological Survey National Geoscience Data 626
Centre (https://www.bgs.ac.uk/services/ngdc/accessions/index.html#item128606).
627
Austrheim, H., Erambert, M., & Boundy, T. M. (1996). Garnets recording deep crustal earthquakes.
628
Earth and Planetary Science Letters, 139(1–2), 223–238. https://doi.org/10.1016/0012- 629
821X(95)00232-2 630
Austrheim, H. (2013). Fluid and deformation induced metamorphic processes around Moho beneath 631
continent collision zones: Examples from the exposed root zone of the Caledonian mountain 632
belt, W-Norway. Tectonophysics, 609, 620–635, 633
https://doi.org/http://dx.doi.org/10.1016/j.tecto.2013.08.030.
634
Behr, W.M., & Platt, J.P., (2014). Brittle faults are weak, yet the ductile middle crust is strong:
635
Implications for lithospheric mechanics. Geophys. Res. Lett. 41, 2014GL061349.
636
https://doi.org/10.1002/2014gl061349 637
Bürgmann, R., & Dresen, G., (2008). Rheology of the Lower Crust and Upper Mantle: Evidence from 638
Rock Mechanics, Geodesy, and Field Observations. Annu. Rev. Earth Planet. Sci. 36, 531–567.
639
https://doi.org/doi:10.1146/annurev.earth.36.031207.124326 640
Carslaw, H.S., & Jaeger, J.C., (1959). Conduction of Heat in Solids. Oxford University Press, Oxford, 641
UK.
642
Chopra, P.N., (1997). High-temperature transient creep in olivine rocks. Tectonophysics 279, 93–111.
643
https://doi.org/10.1016/S0040-1951(97)00134-0 644
Corfu, F., (2004). U–Pb Age, Setting and Tectonic Significance of the Anorthosite–Mangerite–
645
Charnockite–Granite Suite, Lofoten–Vesterålen, Norway. J. Petrol. 45, 1799–1819.
646
Cross, A.J., Prior, D.J., Stipp, M., & Kidder, S., (2017). The Recrystallized Grain Size Piezometer for 647
Quartz: An EBSD-based calibration. Geophys. Res. Lett. 44, 6667–6674.
648
https://doi.org/10.1002/2017GL073836 649
Cross, A. J., & Skemer, P. (2017). Ultramylonite generation via phase mixing in high strain 650
experiments. Journal of Geophysical Research: Solid Earth, n/a-n/a.
651
https://doi.org/10.1002/2016JB013801 652
Dimanov, A., & Dresen, G., (2005). Rheology of synthetic anorthite-diopside aggregates: Implications 653
for ductile shear zones. J. Geophys. Res. Solid Earth 110, 1–24.
654
https://doi.org/10.1029/2004JB003431 655
Di Toro, G., & Pennacchioni, G., (2004). Superheated friction-induced melts in zoned 656
